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2015-11-24
Isotopic signatures of production and uptake of H2 by soil
Chen, Q
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http://dx.doi.org/10.5194/acp-15-13003-2015
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Isotopic signatures of production and uptake of H 2 by soil
Q. Chen1,2, M. E. Popa1, A. M. Batenburg1,3, and T. Röckmann1
1Institute for Marine and Atmospheric research Utrecht, Utrecht University, Utrecht, the Netherlands
2Department of Atmospheric Sciences, University of Washington, Seattle, Washington, USA
3Department of Applied Physics, University of Eastern Finland, Kuopio, Finland Correspondence to: Q. Chen (chenqjie@uw.edu)
Received: 17 April 2015 – Published in Atmos. Chem. Phys. Discuss.: 1 September 2015 Revised: 11 November 2015 – Accepted: 12 November 2015 – Published: 24 November 2015
Abstract. Molecular hydrogen (H2)is the second most abun- dant reduced trace gas (after methane) in the atmosphere, but its biogeochemical cycle is not well understood. Our study focuses on the soil production and uptake of H2 and the associated isotope effects. Air samples from a grass field and a forest site in the Netherlands were collected using soil chambers. The results show that uptake and emission of H2 occurred simultaneously at all sampling sites, with strongest emission at the grassland sites where clover (N2
fixing legume) was present. The H2mole fraction and deu- terium content were measured in the laboratory to deter- mine the isotopic fractionation factor during H2soil uptake (αsoil) and the isotopic signature of H2 that is simultane- ously emitted from the soil (δDsoil). By considering all net- uptake experiments, an overall fractionation factor for de- position ofαsoil=kHD/ kHH=0.945±0.004 (95 % CI) was obtained. The difference in meanαsoilbetween the forest soil 0.937±0.008 and the grassland 0.951±0.026 is not statis- tically significant. For two experiments, the removal of soil cover increased the deposition velocity (vd)andαsoil simul- taneously, but a general positive correlation betweenvdand αsoilwas not found in this study. When the data are evaluated with a model of simultaneous production and uptake, the iso- topic composition of H2that is emitted at the grassland site is calculated asδDsoil=(−530±40) ‰. This is less deuterium depleted than what is expected from isotope equilibrium be- tween H2O and H2.
1 Introduction
H2is considered an alternative energy carrier to replace fos- sil fuels in the future. However, the environmental and cli- mate impact of a potential widespread use of H2is still under assessment. Several studies suggested that the atmospheric H2 mole fraction might increase substantially in the future due to the leakage during production, storage, transportation and use of H2, which could significantly affect atmospheric chemistry (Schultz et al., 2003; Tromp et al., 2003; Van Rui- jven et al., 2011; Warwick et al., 2004).
In the troposphere, H2 has a mole fraction of about 550 parts per billion (ppb=nmol mol−1)and a lifetime of around 2 years (Novelli et al., 1999; Price et al., 2007; Xiao et al., 2007; Pieterse et al., 2011; 2013). H2can affect atmo- spheric chemistry and composition in several ways. Firstly, it increases the lifetime of the greenhouse gas methane (CH4) via its competing reaction with the hydroxyl radical (OH) (Schultz et al., 2003; Warwick et al., 2004). Additionally, H2
affects air quality because it is an ozone (O3)precursor and indirectly increases the lifetime of the air pollutant carbon monoxide (CO) through competition for OH. In the strato- sphere, H2O that is produced through the oxidation of H2
increases humidity, which can result in increased formation of polar stratospheric clouds and O3depletion (Tromp et al., 2003), but this effect may be weaker than estimated initially (Warwick et al., 2004; Vogel et al., 2012).
The main sources of tropospheric H2 are the oxidation of CH4 and non-methane hydrocarbons (NMHC) (48 %), biomass burning (19 %), fossil fuel combustion (22 %) and biogenic N2 fixation in the ocean (6 %) and on land (4 %), while the main sinks are soil uptake (70 %) and oxidation by OH (30 %) (Pieterse et al., 2013).
The biogenic soil sink of H2is the largest and most uncer- tain term in the global atmospheric H2budget. Conrad and Seiler (1981) assumed that the soil uptake of atmospheric H2 is most likely due to consumption by abiotic enzymes, since there were no soil microorganisms known to be able to fix H2 at the low atmospheric mole fraction at that time.
This remained the basic hypothesis of many further soil up- take studies (Conrad et al., 1983; Conrad and Seiler, 1985;
Ehhalt and Rohrer, 2011; Guo and Conrad, 2008; Häring et al., 1994; Smith-Downey et al., 2006). However, Constant et al. (2008a) were first to identify an aerobic microorgan- ism (Streptomyces sp. PCB7) that can consume H2 at tro- pospheric ambient mole fractions and suggested that active metabolic cells could be responsible for the soil uptake of H2rather than extracellular enzymes. Further studies showed that uptake activity at ambient H2level is widespread among the streptomycetes (Constant et al., 2010) and it was postu- lated that high-affinity H2-oxidizing bacteria are the main bi- ological agent responsible for the soil uptake of atmospheric H2 (Constant et al., 2011). Khdhiri et al. (2015) suggested that the relative abundance of high-affinity H2-oxidation bac- teria and soil carbon content could be used as predictive pa- rameters for the H2-oxidation rate. Determining the domi- nant mechanism of the H2soil uptake activity is still an active area of research.
It has been shown that soil uptake of H2can coexist with soil production (Conrad, 1994). H2 is produced in the soil during N2 fixation (e.g., by bacteria living symbiotically in the roots of legumes such as clover or beans) and dark fer- mentation. Although the H2produced in the soil by, e.g., N2
fixation can be largely consumed within the soil, a significant amount of H2escapes to the atmosphere (Conrad and Seiler, 1979, 1980). Conrad and Seiler (1980) estimated that 2.4 to 4.9 Tg a−1of H2is emitted into the atmosphere through N2 fixation on land.
One approach to better understand the sources and sinks of H2 is to investigate the isotopic fractionation processes involved, which act as a fingerprint for H2emitted from dif- ferent sources or destroyed by different sinks. The isotopic composition of H2is expressed as
δ(D,H2)= Rsa
RVSMOW
−1,
where Rsa is the D/H ratio of the sample H2
and RVSMOW=(155.76±0.8) parts per million
(ppm=mmol mol−1) is the same ratio of the standard material, Vienna Standard Mean Ocean Water (VSMOW) (De Wit et al., 1980; Gonfiantini et al., 1993). For brevity, we will use the notation δD (=δD(D, H2))throughout the rest of this paper. TheδD values are usually given in per mill (‰). Recent studies showed that the global meanδD value of atmospheric H2is about+130 ‰ (Batenburg et al., 2011;
Gerst and Quay, 2000, 2001; Rice et al., 2010).
The HH molecule is consumed preferentially over HD dur- ing both OH oxidation and soil uptake, with OH oxidation
causing a much stronger isotope fractionation effect. Only a few studies have investigated the soil uptake of H2 with isotope techniques. Gerst and Quay (2001) carried out field experiments in Seattle, USA, and foundαsoil(=kHD/ kHH) to be 0.943±0.024 (1σ). Note that kHD and kHH are re- moval rate constants for HD and HH, respectively. Rahn et al. (2002a) collected air samples from four forest sites in ecosystems of different ages in Alaska, USA, in July 2001 and obtained a similar average value (0.94±0.01). They sug- gested thatαsoildepends on the forest maturity, with smaller fractionation for more mature forests. Since the more ma- ture forests showed larger deposition velocity (vd) of H2, they further suggested that lower uptake rates involve greater isotopic fractionation (αsoil further from 1) than fast uptake rates. Rice et al. (2011) performed deposition experiments in Seattle and foundαsoil varying from 0.891 to 0.976, with a mean of 0.934. They foundαsoilto be correlated withvd, with smaller isotope effects (αsoilcloser to 1) occurring at higher vd, which agreed with the suggestion by Rahn et al. (2002a).
In addition, unpublished experiments from Rahn et al. (2005) yieldedαsoil=0.89±0.03 in three upland ecosystems that were part of an Alaskan fire chronosequence. The data sug- gest that variability in the soil/ecosystem affectsαsoilbut no significant variability ofαsoilwith season was detected. Hith- erto, onlyαsoilvalues from studies in Seattle and Alaska are available, and values from other locations and ecosystems are needed to learn more about the factors influencingαsoil.
The δD of H2 from various surface sources has been reported as about −290 ‰ for biomass burning (Gerst and Quay, 2001; Haumann et al., 2013) and between
−360 and −200 ‰ for fossil fuels combustion (Rahn et al., 2002b; Vollmer et al., 2012). So far no field stud- ies have determined the isotopic composition of H2 emit- ted from soil. Two laboratory studies examined the iso- topic signature of H2 produced from N2 fixation. Luo et al. (1991) reported a fractionation factorαH2/H2O=R(D/H, H2) / R(D/H, H2O)=0.448±0.001 between the H2 pro- duced from N2 fixation and the H2O used to grow the N2- fixing bacteria for Synechococcus sp. and 0.401±0.002 for Anabaena sp., respectively. Walter et al. (2012) reported αH2/H2O=0.363±0.019 for the N2-fixing rhizobacterium Azospirillum brasiliensis. It has been proposed that micro- biological H2consumption and production could modify the thermal isotopic equilibrium between H2 and H2O in low- temperature hydrothermal fluids (Kawagucci et al., 2010).
Compared to the surface sources, H2produced from CH4and NMHC oxidation is isotopically strongly enriched in deu- terium, withδD between +120 and +180 ‰ (Rahn et al., 2003; Röckmann et al., 2003a; Pieterse et al., 2011).
Here we report measurements of the isotopic fractionation factors of H2during soil deposition at two different sites in the Netherlands, a forest and a grassland site. For the grass- land site we also determine the apparent isotopic composition of H2that was simultaneously emitted from soil during the experiment.
2 Methods 2.1 Sampling
Air samples were collected from a soil chamber at two lo- cations in the Netherlands (Fig. 1): a grass field around the Cabauw tall tower (51◦580N, 4◦550E) and a forest site near Speuld (52◦130N, 5◦390E). Two types of ground cover (grass with and without clover) were sampled at Cabauw, while three types of forest (Douglas fir, beech and spruce) were se- lected in Speuld. More information about the soil and vegeta- tion type can be found in Beljaars and Bosveld (1997) for the Cabauw site and in Heij and Erisman (1997) for the Speuld site.
Flask samples were filled with air from a soil chamber, using a closed-cycle air sampler (Fig. 2). The soil cham- ber consisted of two parts: the chamber body with a metal base at the bottom that was inserted about 2 cm into the soil and a removable transparent lid with two connections for air sampling. The chamber had a height of 40 cm, an area of 570 cm2and a volume of 22.8 L; the air inside was mixed by a fan. The sampler could hold four flasks installed in series, which could be bypassed independently; the flow and pres- sure in the flasks were controlled. The air was dried using Mg(ClO4)2. After passing through the flasks the air was re- turned to the soil chamber, which kept the pressure inside the chamber approximately constant during sampling.
Air samples were collected from the chamber in 1 L glass flasks at 0, 10, 20 and 30 min after closing the chamber (time interval changed to 5 min in Speuld because of the faster uptake). The gas flasks (Normag, Ilmenau, Germany) were made of borosilicate glass 3.3 with O-ring-sealed stop- cocks made of PCTFE (Kel-F) and covered with a dark hose.
Thorough tests have demonstrated that air samples with typ- ical trace gas content are stable in these flasks (Rothe et al., 2004). In the beginning, the whole sampling unit (all lines, connections and flasks) was flushed with ambient air for about 10 min at a flow rate of 2 L min−1and a pressure of 100 kPa, with all flasks open and the chamber lid open.
This initial flushing process was designed to fill the flasks with background air. The air pressure inside the flasks was increased to 200 kPa (180 kPa for Speuld samples) by adjust- ing the flow control valve and the valves on two pressure gauges (Fig. 2) before chamber closing and then maintained constant during the whole sampling time. The flow rate was maintained at 2 L min−1 at ambient pressure and tempera- ture with a rotameter and the pressure inside the chamber was maintained at 100 kPa during the whole sampling time.
The temperature was not recorded during the sampling. Af- ter the initial flushing, the first flask was closed and then the chamber was closed as well. Afterwards, the air was flushed from the chamber through three flasks (the first flask was by- passed) and back to the chamber. After 10, 20 and 30 min, the second, third and fourth flasks were closed.
A total of 36 sets of air samples were collected in Cabauw during summer (June, July and August) 2012 and 12 sets were collected in Speuld in September 2012. Each set con- tains four air samples. In total, 186 valid samples were an- alyzed for H2 mole fraction and its deuterium content (six were lost during sampling, transportation and measurement).
All the Speuld samples and about half of the Cabauw sam- ples were further used for analysis in this study. The reason why 50 % of the Cabauw experiments were not used is that these experiments showed neither strong H2emission nor H2 uptake and the isotopic signals were weak. Most experiments were conducted with the 22.8 L volume soil chamber as de- scribed above, while 10 experiments were conducted with a larger automated soil chamber with a volume of 125 L and a height of 22.5 cm.
2.2 Laboratory determination of H2mole fraction and deuterium content of air samples
The mole fraction and δD of H2 were measured with a gas chromatography isotope ratio mass spectrometry (GC/IRMS) setup (Rhee et al., 2004). For H2 mole frac- tions, the laboratory working standards are linked to the MPI- 2009 scale (Jordan and Steinberg, 2011). TheδD values of the laboratory reference gases are indirectly linked to mix- tures of synthetic air with H2 of known isotopic composi- tion, certified by Messer Griesheim, Germany (Batenburg et al., 2011). Most of the samples collected from Cabauw were measured within 2 months after sampling, while the samples from Speuld were kept in a dark storage room for around 4 months before measurement.
The operational principle of the GC/IRMS system is to separate H2 from the air matrix at low temperature (about 36 K) and measure the HH and HD content with a mass spec- trometer. The measurement includes four main steps.
– A glass sample volume (750 mL) is evacuated and subsequently filled with sample air to approximately 700 mbar. This volume is then exposed to a cold head (36 K) of a closed-cycle helium compressor for 9 min.
During this stage, all gases except H2, helium (He) and neon (Ne) condense.
– The remainder in the headspace of the cold head and sample volume is then flushed with He carrier gas to a pre-concentration trap where H2is collected on a 25 cm long, 1/8 inch OD (outside diameter) stainless steel tube filled with fine grains (0.2 to 0.5 mm) of 5 Å molecular sieve, for 20 min. The pre-concentration trap is cooled down to the triple point of nitrogen (63 K) by keeping it in a liquid N2reservoir that is further cooled down by pumping on the gas phase.
– After the collection of H2, the pre-concentration trap is warmed up to release the absorbed H2, which is then cryo-focused for 4 min on a capillary (25 cm long,
Beech Grass
Douglas fir
Spruce
Clover
Cabauw Speuld
Netherlands
Figure 1. The location of the two sampling sites (Cabauw and Speuld) in the Netherlands, as well as the plant species there.
Chamber
Flask 1 Flask 2 Flask 3 Flask 4
Mg(ClO4)2 Pump Filter
Rotameter Pressure gauge
Pressure gauge
Fan Flow control valve
Figure 2. Scheme of the sampling setup using the closed-cycle air sampler. The volume of the soil chamber was 22.8 L and the volume of each flask was 1 L.
0.32 mm inside diameter) filled with 5 Å molecular sieve at 77 K. After that, the cryo-focus trap is warmed up to ambient temperature and the H2sample is flushed with He carrier gas onto the GC column (5 Å molec- ular sieve, ≈323 K) where H2 is chromatographically purified from potential remaining interferences.
– In the end, the purified H2is carried by the He carrier gas via an open split interface (Röckmann et al., 2003b) into the IRMS for D/H ratio determination.
More details about the GC/IRMS system and measure- ment steps can be found in Rhee et al. (2004) and Röck- mann et al. (2010). The data correction procedures and iso- tope calibration are similar to those described in Baten- burg et al. (2011). Four reference gases were used to de- termine the δD values of the samples. Two of them (Ref- 1 and Ref-2) with δD values of (+207.0±0.3) ‰ and (+198.2±0.5) ‰ were calibrated and used previously in Batenburg et al. (2011). The other two new reference gases (Ref-3 and Ref-4) were calibrated vs. Ref-1 and Ref-2. The δD value of Ref-3 was (−183±2.4) ‰. Ref-4 was a fre-
quently measured reference gas that was measured usually about five times per sequence of measurement, while other three reference gases were measured about one to three times per sequence of measurement. TheδD value of Ref- 4 dropped linearly with time from−115 to−157 ‰ between 1 June 2012 and 15 February 2013, while the other three ref- erence gases were stable.
2.3 Non-linearity of the GC/IRMS system
Ideally, theδD of H2measured with the GC/IRMS should not depend on the total amount of H2used for analysis, but in practice a dependence of the isotopic composition on the amount of H2 is observed for low mole fractions. This is called non-linear behavior, and it is a particularly severe lim- itation for soil uptake studies, since the mole fraction in such samples can decrease by more than an order of magnitude.
For comparison, in ambient background air the H2mole frac- tion variations are usually no more than 20 %.
Experiments were carried out with different quantities of air from various laboratory reference bottles with knownδD
−2 −1 0 1 2 3
−150
−100
−50 0 50
ln (peak area)
δD difference (‰)
Ref1 Ref2 Ref3 S1 S2 S3 S4 S5 S6 S7 Linear 95% CI
Figure 3. Difference of δD from the assigned value for different gases including reference gases (Ref1-3) and laboratory flask sam- ples (S1-7). A linear function (y=54.6x) was fit to the data with peak area between 0.2 and 1.0 Vs (green solid line; the dashed lines represent the 95 % confidence interval of the fit). This function was used to correct the soil experiment data that were measured at low peak areas.
to determine a suitable correction for the non-linear behav- ior. The measuredδD increases with the mass 2 sample peak area, which is proportional to the H2quantity in the sample.
In the peak area range of 0.2 Vs to 1 Vs this relation can be parameterized by a logarithmic functionδD=54.6 ln (peak area (Vs)−1) ‰, which is used as correction function for the measurements at low peak areas (Fig. 3). The linearity cor- rection introduces an additional uncertainty due to uncertain- ties in the logarithmic fit, particularly at low peak areas. The total assigned uncertainty for each measurement is calculated from the analytical and fitting uncertainty, as a function of peak area (Fig. 4). It is 2 ‰ for ln (peak area (Vs)−1) of 1.5 or more (equivalent to more than 600 ppb H2in an air sam- ple) but increases to 32 ‰ when ln (peak area (Vs)−1) drops to−1.6 (≈20 ppb H2in air sample). In total, theδD results of 18 Speuld samples that were measured at these low peak areas were corrected with this linearity correction. Possible additional systematic errors (a few ‰ ) may arise from uncer- tainties in the initially assignedδD values of the commercial calibration gases, changes of these values in the process of creating calibration mixtures with near-ambient H2concen- tration, and the calibration measurements themselves (Baten- burg et al., 2011).
2.4 Data evaluation
Assuming first order kinetics for H2removal and a constant production ratePover the course of a deposition experiment, the time evolution of the mole fractioncof non-deuterated H2(HH) inside the soil chamber can be expressed as
dc
dt =P−kc, (1)
−2 −1.5 −1 −0.5 0 0.5 1 1.5
−40
−30
−20
−10 0 10 20 30 40
ln (peak area)
δD uncertainty (‰)
Figure 4. Calculated total assigned uncertainty ofδD (consisting of analytical uncertainty and uncertainty arising from the linearity correction) for air samples with ln(peak area) ranging from−1.6 to 1.5.
wherek is the first order uptake rate constant of HH. For well-mixed air in the chamber,k=vd/ h, where vd is the gross deposition velocity of H2andhis the chamber height.
The gross deposition velocity is the deposition velocity cor- rected for production, which is different from the net deposi- tion velocity reported in some studies in the past that showed the effective uptake of H2from the atmosphere. The solution of Eq. (1) is of the form
c=(ci−ce) e−kt+ce, (2)
wherec,ci andce(=P / k)are the mole fractions of HH at timet, initially and at equilibrium, respectively. Therefore, P andkcan be obtained by fitting an exponential function to the time evolution of HH inside the chamber. Similarly, we can obtainP0andk0from the time evolution of HD.
c0= ci0−c0e
e−k0t+c0e, (3)
wherec0,c0i,c0e(=P0/ k0),P0 andk0are the corresponding parameters for HD.
Equations (2) and (3) constitute the mass balance model that we used to analyze our data. Whenk,k0,P andP0have been determined,αsoilandδDsoilcan be calculated simply as αsoil=k0
k (4)
δDsoil= P0/P 2RVSMOW
−1. (5)
However, fitting an exponential curve to only four sample data yields relatively large errors fork,k0,PandP0, which propagate to large errors forαsoilandδDsoilif they are deter- mined directly from Eqs. (4) and (5).
In Rice et al. (2011), Eqs. (2) and (3) were combined to calculateαsoil in the presence of both source and sink of H2 usingceandc0efrom the exponential fits:
lnc0−c0e ci0−ce0 =k0
k lnc−ce
ci−ce
. (6)
αsoil=k0/ kcan be obtained by plotting lnc
0−ce0
c0i−c0e vs. lncc−ce
i−ce
and fitting a linear function. In the absence of soil emission (ce=c0e=0), Eq. (6) collapses to the well-known Rayleigh fractionation equation that is used to quantify the isotope fractionation during single stage removal processes in the ab- sence of sources.
For the high emission measurements, where production overwhelms consumption, we use the relations ce=P / k and c0e=P0/ k0 and obtain P0/ P from the slope of c0elnc
0−ce0
c0i−c0e againstcelncc−ce
i−ce. ThenδDsoil is calculated from Eq. (5).
2.5 Flask sampling model
The advantage of sampling with the soil chamber system de- scribed in Sect. 2.1 was that the pressure in the soil chamber stayed constant even when several large samples (2 L each) were taken. A disadvantage was that the volume of air inside the flasks (8 L of air in total) was considerable compared to the volume of air inside the soil chamber (22.8 L). This had two effects: (1) a significant part of the air was at each time separated from the chamber and thus from the soil produc- tion and uptake and, (2) because of the time lag to flush the samples, the air in a flask was not the same as the air in the chamber at the same time.
We built a flask sampling model to derive correction fac- tors that take into account the influence of the flask sampling system. For a given combination of uptake and production rates, the model simulates the evolution of the H2mole frac- tion in two configurations: the soil chamber alone and the soil chamber plus four flasks as in our experiments. The model is described in detail in Appendix A. An example of a simu- lation is shown in Fig. 5. Compared to the situation without flasks, there is a time lag in the decay of H2 for both the chamber and the flasks after introducing four flasks in the model. The time lag for the second flask is about 2.5 min. It increases to 5 min for the third flask and is even longer for the fourth flask.
It is obvious that the sampling process strongly affects the uptake ratekappand production ratePapp obtained from the direct flask measurements, so we corrected allkappandPapp
values with the correction coefficients derived from this flask sampling model (Appendix A). For a fixed chamber volume, sample pressure, flow rate and time interval of the flask col- lection that are all recorded for each experiment, the rela- tionship between the actual uptake rate constant ktrue and apparent uptake rate constantkappcan be obtained (see Ap- pendix A). Under the same sampling conditions for a fixed
0 5 10 15 20 25 30
100 150 200 250 300 350 400 450 500 550
7ime (min) H 2 mole fraction (ppb)
hamber − with flasks flask 2
flask 3 flask 4
hamber − without flasks C
C
Figure 5. Results of the flask sampling model with the fol- lowing parameters:k=0.1 min−1, P=10 ppb min−1and c1(t= 0)=530 ppb. The figure shows the evolution of H2mole fraction in the chamber (green curve), in flask 2 (blue curve), flask 3 (red curve) and flask 4 (magenta curve) as a function of time and what would be expected for a chamber without flasks (black curve). Flask 1 was closed before closing the chamber (at time 0 when all volumes con- tained the same air).
value of Papp, the relationship between actual production ratePtrueand apparent production ratePappdepends onktrue
(Fig. 10b).
To evaluate the data, we first applied an exponential fit as in Eq. (2) to the measured HH mole fractions for the four flasks in each experiment and obtained apparent valueskapp, Papp andce,app from the fit parameters. Then we used the correction factors derived from the flask sampling model to retrieve true valuesktrue andPtruefrom the apparent values kappandPapp. One can obtaink0trueandPtrue0 by applying the same method to HD mole fractions inside four flasks.
To determine αsoil, we plotted lnc
0−c0e,app
c01−c0e,app vs. lncc−ce,app
1−ce,app
(Eq. 6, Fig. 7) and obtainedαsoil,appfrom the slope of the lin- ear regression. Here,candc0are HH and HD mole fractions in each of the four flasks;c1andc01are HH and HD mole frac- tions of the first flask;ce,appandce,app0 are apparent HH and HD equilibrium mole fractions obtained from the exponen- tial fits of HH and HD mole fractions inside the four flasks.
We determined the relationship (Fig. 10c) betweenαsoil,true
andαsoil,app obtained from lnc
0−c0e,app
c01−c0e,app vs. lncc−ce,app
1−ce,app using the flask sampling model (see Appendix A1.3). The correc- tion coefficients for each experiment are given in Table 3.
Similarly, we obtained Papp0 /Papp by plotting ce,app0 lnc
0−c0e,app
c01−c0e,app vs. ce,applncc−ce,app
1−ce,app (Fig. 9), and cal- culated δDsoil,app by use of Eq. (5). Then we retrieved δDsoil,true by use of the flask sampling model (Fig. 10d).
The corresponding correction coefficients forδDsoil,app for each net-emission experiment are shown in Table 3. More
Figure 6. Time evolution of H2, HD andδD in Cabauw (upper and middle panels) and in Speuld (lower panel) for representative experiments.
HD is calculated from H2andδD. The H2data are fitted with an exponential function of the formc= c1−ce,app
e−kappt+ce,app, where c1 andce,app are the H2mole fractions initially and in equilibrium, andkapp is the apparent soil uptake rate constant for H2. A similar exponential function is applied to the HD data. Error estimates for H2, HD andδD are shown. The connecting lines forδD data are included to guide the eye.
information about the retrievals of αsoil,true and δDsoil,true can be found in Appendix A.
Overall, the sampling effect on δDsoil is small (less than 22 ‰). This means that the flask sampling system strongly affects the temporal evolution of HH and HD individually (Fig. 5), and the uptake and production rates derived from flask measurements, but the effects on the computed isotopic signature of the source and sink are relatively small. More de-
tails and discussion of the flask sampling model corrections are provided in Appendix A.
3 Results
3.1 Temporal evolution of H2, HD andδD
Figure 6 shows examples for the temporal evolution of H2, HD andδD in Cabauw and Speuld, with error estimates in-
−8 −6 −4 −2 0
−8
−6
−4
−2 0
ln[(c−ce,app)/(c1−ce,app)]
ln[(c’ −c’ e,app)/(c’ 1−c’ e,app)]
SPU CBW Overall fit
Figure 7. Plot of lnc
0−c0e,app
c01−c0e,app vs. lncc−ce,app
1−ce,app for all Speuld and Cabauw net-uptake experiments. The slope of the linear fit to the data returns the fractionation factorαsoil,app=0.947±0.004 (95 % CI). Errors inxandydirection for each data point were considered.
One outlier (“CBW-18”) was not included in the fitting. The 95 % confidence intervals of the fit line are included as dashed lines but largely overlap with the fit line.
0 0.05 0.10 0.15 0.20 0.25 0.30
0.85 0.90 0.95 1.00 1.05
vd (cm s−1) α soil
SPUCBW
Figure 8. Correlation betweenαsoil and vd for all Speuld exper- iments and Cabauw net-uptake experiments. The errors for αsoil were taken from Table 1.
−10000 −8000 −6000 −4000 −2000 0
−1.6
−1.4
−1.2
−1.0
−0.8
−0.6
−0.4
−0.2 0
ce,appln[(c−ce,app)/(c1−ce,app)]
c’ e,appln[(c’ −c’ e,app)/(c’ 1−c’ e,app)]
CBW−8 CBW−10 CBW−14 CBW−17 CBW−21 CBW−28 CBW−30 CBW−31 CBW−33
Figure 9. Plot ofc0e,applnc
0−ce,app0
c01−ce,app0 vs.ce,applncc−ce,app
1−ce,app for nine Cabauw net-emission experiments. A linear function was fit to each individual data set and the slope was used to calculate theδDsoil,app value for each experiment. Errors inxandydirection for each data point were considered.
cluded. The errors for H2and HD are about 4 % of the re- spective mole fraction. The error for δD ranges from 2 to 17 ‰.
Some of our Cabauw experiments show net soil emission of H2 (upper panels) and some show net soil uptake (mid- dle panels), while all Speuld experiments show net uptake of H2(lower panels). In the Cabauw net-emission experiments, the increase in H2mole fractions is associated with a strong decrease inδD, showing a strongly depleted H2source. How- ever, the net-uptake experiments at Cabauw show also a de- crease inδD, albeit smaller. In the Speuld experiments, the uptake of H2is much faster; theδD increases in the begin- ning but then decreases again towards the end of the sam- pling, when the H2mole fractions are low.
As mentioned in the introduction, soil uptake tends to in- creaseδD while soil emission tends to decreaseδD of H2. The continuous decrease ofδD with time in all Cabauw ex- periments and the eventual decrease ofδD in all Speuld ex- periments clearly show that there is concurrent soil emission even with net uptake. Thus, the equilibrium H2concentration in our experiments is not just a threshold concentration where microbial uptake stops, but the isotopic evolution shows that there is an active overlapping emission (Conrad, 1994).
Table 1. The deposition velocity (vd), fractionation factor (αsoil)as well as its error estimate and soil cover information for each Speuld experiment (a) and Cabauw net-uptake experiment (b). The SD represents standard deviation and SE represents standard error. The errors of αsoilrepresent the 95 % confidence interval (CI) forαsoil,appobtained from lnc
0−c0e,app
c10−ce,app0 vs. lncc−ce,app
1−ce,app. (a) Fn(nmol m−2s−1) vd(cm s−1) αsoil Errorαsoil Soil cover
SPU-1 −30.1 0.20 0.924 0.032 D. fir, moss
SPU-2 −35.3 0.22 0.948 0.028 D. fir, needles
SPU-3 −37.7 0.20 0.945 0.008 D. fir, moss
SPU-4 −26.1 0.16 0.913 0.004 D. fir, moss
SPU-5 −24.9 0.16 0.918 0.006 D. fir, moss
SPU-6 −13.2 0.12 0.951 0.031 D. fir, moss
SPU-7 −19.6 0.12 0.939 0.005 beech, leaves
SPU-8 −28.4 0.16 0.955 0.008 same subsite as SPU-7, leaves removed
SPU-9 −20.4 0.12 0.925 0.002 beech, leaves
SPU-10 −22.3 0.13 0.949 0.060 spruce, moss
SPU-11 −19.4 0.13 0.936 0.068 spruce, needles
SPU-12 −40.5 0.28 0.947 0.004 same subsite as SPU-11, needles removed
MEAN −26.5 0.17 0.937 – –
SD 8.2 0.05 0.014 – –
SE 2.4 0.01 0.004 – –
(b) Fn vd αsoil Errorαsoil Soil cover
(nmol m−2s−1) (cm s−1)
CBW-5 −6.6 0.04 0.943 0.004 few clover, grass
CBW-7 −3.1 0.03 1.019 0.005 few clover, grass
CBW-16 −22.9 0.18 0.993 0.001 bare soil, few grass
CBW-18 −39.3 0.24 0.950 0.054 grass
CBW-19 −7.4 0.14 0.935 0.105 grass
CBW-20 −14.9 0.20 0.940 0.260 bare soil
CBW-25 −8.0 0.12 0.911 0.014 clover, grass
CBW-26 −6.1 0.09 0.916 0.038 grass
MEAN −13.6 0.13 0.951 – –
SD 12.2 0.08 0.037 – –
SE 4.3 0.03 0.013 – –
3.2 Emission and uptake strength of H2
The production rate P =Ptrue and uptake rate constant k=ktrue were obtained by applying exponential fits to the temporal evolution of H2 and applying the corrections de- rived from the flask sampling model (Appendix A) to the Pappandkappobtained from the exponential fits (Fig. 6). The deposition velocity (vd), production flux (Fp), initial uptake flux (Fu)and net flux at the beginning of the experiment (Fn) were then calculated as follows:
vd=kh, (7)
Fp=P h
VM, (8)
Fu=kc1h VM
, (9)
Fn=Fp−Fu, (10)
whereh,VMandc1are the chamber height, standard molar volume (=22.4 L mol−1)and H2 mole fraction of the first flask, respectively. We note that with our method we derive vd as deposition velocity for the gross uptake, unlike most of the results reported in the literature that just measured net uptake.
The strongest soil uptake occurs in the Speuld experi- ments (Table 1a), with a mean vd of (0.17±0.02) (2 SE, n=12) cm s−1 (SE represents standard error). On average, the Cabauw experiments show weaker soil uptake, with a mean vd of (0.13±0.06) (2 SE, n=8) cm s−1 for the net-uptake experiments (Table 1b) and (0.06±0.03) (2 SE, n=9) cm s−1 for the net-emission experiments (Table 2).
In terms of the net H2fluxFn, this is (−26.5±4.8) (2 SE, n=12) nmol m−2s−1 for Speuld experiments (Table 1a), (−13.6±8.6) (2 SE, n=8) nmol m−2s−1 for Cabauw net- uptake experiments (Table 1b) and (49.5±29.8) (2 SE, n=9) nmol m−2s−1for Cabauw net-emission experiments
0 0.05 0.10 0.15 0.20 0.25 0.30 1.40
1.45 1.50 1.55 1.60 1.65
kapp (min−1) k true/k app
Ptrue=50 ppb min−1 Ptrue=200 ppb min−1 Ptrue=650 ppb min−1
0 50 100 150 200 250 300 350 400 450 1.40
1.45 1.50 1.55 1.60
Papp (ppb min−1) P true/P app
ktrue=0.05 min−1 ktrue=0.25 min−1 ktrue=0.45 min−1
0.90 0.92 0.94 0.96 0.98 1.00
0.98 0.99 1.00
αsoil,app α soil,true/α soil,app
(δDsoil,true+1)=0.25 (δDsoil,true+1)=0.45 (δDsoil,true+1)=0.65
0.25 0.3 0.35 0.4 0.45 0.5 0.55 0.6 0.65 0.99
1.00 1.01 1.02 1.03 1.04 1.05
(δDapp+1) (δD true+1)/(δD app+1)
αsoil,true=0.90 αsoil,true=0.95 αsoil,true=1.00
(a) (b)
(d) (c)
Figure 10. (a) The relationship betweenktrue/ kappandkappforPtrueof 50, 200 and 650 ppb min−1; (b) betweenPtrue/ PappandPappfor ktrueof 0.05, 0.25 and 0.45 min−1; (c) betweenαsoil,true/ αsoil,appandαsoil,appfor (δDsoil,true+1) of 0.25 to 0.65 forktrue=0.25 min−1and Ptrue=50 ppb min−1; (d) between (δDsoil,true+1)/(δDsoil,app+1) and (δDsoil,app+1) forαsoil,trueof 0.90 to 1.00 forktrue=0.25 min−1 andPtrue=50 ppb min−1. The parameters of the sampling setup areV0=22.8 L,f =2 L min−1,1t=10 min and the pressures inside the flasks and chamber are 200 and 100 kPa, respectively.
(Table 2), indicating strong uptake, weaker uptake and strong emission of H2, respectively.
3.3 Fractionation during soil uptake
Soil uptake and soil emission have opposite effects on the isotopic composition of H2and can partly cancel each other.
This will lead to additional uncertainty and we expect to ob- tain the most robust fractionation factor for soil uptake when the soil uptake is larger than the soil emission (Table 1a, b).
The resultingαsoilfor Speuld (Table 1a) varies from 0.913 to 0.955, with a mean value of 0.937±0.008 (2 SE,n=12).
Error estimates for HH and HD mole fraction at time t and at equilibrium are considered for the final error estimates of αsoilfor each experiment.
Table 1b shows αsoil of the Cabauw net-uptake experi- ments. It should be noted that the soil-emitted H2interferes
much more with the fractionation during uptake in these Cabauw net-uptake experiments than in the Speuld experi- ments, which is illustrated by the consistent decrease inδD in the middle panel of Fig. 6. The derived values forαsoil
vary from 0.911 to 1.019 with a mean value of 0.951±0.026 (2 SE,n=8) for these eight selected Cabauw net-uptake ex- periments. Both the mean and the standard error are higher than in the Speuld experiments (0.937±0.008), but the dif- ference is not significant at the 0.1 confidence level.
To graphically illustrate the calculation ofαsoil with the mass balance model, we plot lnc
0−c0e,app
c01−c0e,app vs. lncc−ce,app
1−ce,app for all Speuld and Cabauw net-uptake experiments in Fig. 7. A linear fit is applied to all the data and the overallαsoil,app is found to be 0.947±0.004 (95 % CI). Applying a correction factor is not straightforward now because this analysis com- bines the results from different experiments. If we use the
Table 2. Net flux, deposition velocity andδDsoil (including error) obtained from the mass balance model for the net H2emission ex- periments.
Net emission Fn vd δDsoil ErrorδDsoil
(nmol m−2s−1) (cm s−1) (‰) (‰)
CBW-8 24.5 0.05 −535 53
CBW-10 16.1 0.03 −460 17
CBW-14 13.7 0.02 −629 21
CBW-17 20.3 0.03 −542 1
CBW-21 42.0 0.04 −574 3
CBW-28 150.2 0.14 −488 83
CBW-30 41.0 0.05 −580 7
CBW-31 92.0 0.09 −509 7
CBW-33 46.2 0.10 −451 52
MEAN 49.5 0.06 −530 –
SD 44.7 0.04 59 –
SE 14.9 0.01 20 –
average ofαsoil,true/ αsoil,appratios (0.998) for all net-uptake experiments in Table 3 as the correction coefficient for this overallαsoil,app, the overallαsoilis 0.945±0.004 (95 % CI).
Figure 8 shows αsoil as a function of vd for all Speuld experiments and Cabauw net-uptake experiments. The R2 value is nearly 0 and the p value is 0.53 for the linear re- gression of all experiments, so no significant correlation be- tweenαsoilandvdis found. Also, no significant correlation is found when considering the Speuld and Cabauw net-uptake experiments separately.
3.4 Isotopic signature of H2emitted from soil
As discussed in Sect. 2.4, the isotopic signature of H2emit- ted from the soil (δDsoil)can be obtained from the mass bal- ance model. In order to minimize the influence of soil uptake on the computedδDsoiland obtain the most robust result, we only consider the Cabauw experiments with strong soil emis- sion and weak soil uptake (ce,app> 1500 ppb). In total, nine Cabauw experiments are selected (Table 2) and a linear fit is applied to the plot ofc0e,applnc
0−c0e,app
c10−c0e,app vs.ce,applncc−ce,app
1−ce,appfor each experiment (Fig. 9). It can be seen that the linear func- tion fits the data very well for each experiment. The slope of the linear fit yields Papp0 / Papp. ThisPapp0 / Papp ratio is used to calculate δDsoil,app (Eq. 5). After correcting for the flask sampling effects (see Appendix A), the corresponding δDsoil values are shown in Table 2. TheδDsoil value ranges from−629 to−451 ‰, with a mean value of (−530±40) ‰ (2 SE,n=9), which is very D depleted, but still considerably enriched relative to the value around −700 ‰ expected for thermodynamic equilibrium between H2and H2O (Bottinga, 1969).
4 Discussion
4.1 Emission and uptake strength of H2
The deposition velocityvd is a measure of the strength of soil uptake. Both microbial removal and diffusion can affect vd, and they can both be influenced by the temperature and moisture content of the soil (Ehhalt and Rohrer, 2013a, b).
On average, thevdobtained in this study is larger in the forest region (Table 1a) than in the grass/clover region (Table 1b and 2), in agreement with the conclusion from Ehhalt and Rohrer (2009).
Thevdof (0.06±0.03) cm s−1found in our Cabauw net- emission experiments (Table 2) is similar to those reported in Conrad and Seiler (1980) (0.07 cm s−1, both grass and clover) and Gerst and Quay (2001) (0.04 cm s−1, grass), while thevd of (0.13±0.06) cm s−1 in Cabauw net-uptake experiments (Table 1b) is larger than those studies with simi- lar soil cover but close to values of 0.12 to 0.14 cm s−1found in savanna soil (Conrad and Seiler, 1985). The stronger soil uptake in Speuld forest ((0.17±0.02) cm s−1) agrees well with the beech forest results (0.06 to 0.22 cm s−1)in Förs- tel (1988) and Förstel and Führ (1992). However, other stud- ies at forest sites cited in Ehhalt and Rohrer (2009) showed lower vd than our Speuld results. We note here that the vdvalues reported in Conrad and Seiler (1980, 1985) were gross deposition velocities while those reported in Gerst and Quay (2001) were net deposition velocities. The specific method used to obtainvd was not documented in the other studies.vd values obtained from our experiments are gross deposition velocities.
The net-uptake flux Fn in our Speuld experiments and Cabauw net-uptake experiments is much larger than those found in Smith-Downey et al. (2008). They found aFn of about −8 nmol m−2s−1 for the forest, desert and marsh, which was similar to that for loess loamy soil in Schmitt et al. (2009). Our results are within the Fn range found in the mixed wood plains by Constant et al. (2008b) and the Harvard forest by Meredith (2012). Previously at our Cabauw site, Popa et al. (2011) obtained a Fn of only
−3 nmol m−2s−1 by using the radon tracer method. How- ever, the Cabauw net-uptake experiments used for this eval- uation were from selected places where uptake was strong, while the results in Popa et al. (2011) represented the overall uptake in the footprint of the Cabauw site, which is a much larger area (tens of km2).
Khdhiri et al. (2015) performed microbiological analyses on soil samples from the Cabauw and Speuld sites in order to find the drivers of soil H2uptake. They observed that the H2 uptake rate under standard incubation conditions was signifi- cantly lower for the Cabauw soil samples than for the Speuld ones, which is consistent with our findings. The main factors that explained the differences were the relative abundance of high-affinity H2-oxidizing bacteria and the soil carbon con- tent, both lower on average for the Cabauw site.
Table 3. Sampling information and the correction coefficients (ktrue/ kapp, Ptrue/ Papp, αsoil,true/ αsoil,app and (δDsoil,true+1)/ (δDsoil,app+1) used for each experiment. Size S refers to small chamber and size L refers to large chamber.
Exp. Pressure Flow rate Size 1t kapp Papp ktrue/ kapp Ptrue/ Papp αsoil,true/αsoil,app (δDsoil,true+1)/
(kPa) (L min−1) (min) (min−1) (ppb min−1) (δDsoil,app+1)
SPU-1 200 2 S 10 0.199 4.12 1.494 1.601 0.984 –
SPU-2 200 2.2 S 5 0.206 0.67 1.589 7.472 0.998 –
SPU-3 200 3.1 S 5 0.204 3.58 1.496 2.475 0.999 –
SPU-4 200 2.8 S 5 0.160 7.51 1.526 2.136 1.004 –
SPU-5 200 2.6 S 5 0.156 4.16 1.546 2.759 1.004 –
SPU-6 160 3.2 L 5 0.232 7.61 1.184 1.446 0.999 –
SPU-7 160 3.2 S 5 0.128 5.40 1.418 2.264 1.006 –
SPU-8 160 2.5 S 5 0.172 4.23 1.438 2.381 1.001 –
SPU-9 160 2.8 S 5 0.128 4.56 1.440 2.513 1.007 –
SPU-10 180 2.7 S 5 0.128 – 1.502 / 1.005 –
SPU-11 160 2.2 S 5 0.130 – 1.490 / 1.006 –
SPU-12 180 2.3 S 5 0.272 11.30 1.529 1.720 0.994 –
CBW-5 200 2 L 10 0.086 18.24 1.204 1.248 1.001 –
CBW-7 200 1.9 L 10 0.048 11.57 1.260 1.361 0.999 –
CBW-16 210 2.1 S 10 0.183 45.21 1.498 1.505 0.999 –
CBW-18 200 2 S 10 0.240 38.07 1.532 1.527 0.986 –
CBW-19 200 2 S 10 0.145 56.69 1.457 1.463 0.991 –
CBW-20 200 2 S 10 0.196 65.81 1.491 1.494 0.988 –
CBW-25 200 2 S 10 0.122 44.85 1.449 1.460 0.994 –
CBW-26 200 2 S 10 0.088 31.05 1.452 1.475 1.002 –
CBW-8 200 2 S 10 0.044 82.92 1.542 1.438 – 1.048
CBW-10 200 2.6 L 10 0.069 111.00 1.177 1.152 – 1.010
CBW-14 200 2.5 L 10 0.035 82.53 1.251 1.166 – 1.042
CBW-17 220 2.1 L 10 0.047 117.40 1.268 1.198 – 1.024
CBW-21 220 2 L 10 0.078 232.20 1.209 1.179 – 1.008
CBW-28 175 1.8 S 10 0.146 440.90 1.412 1.402 – 1.018
CBW-30 200 2 L 10 0.090 237.70 1.202 1.180 – 1.008
CBW-31 200 2 S 10 0.098 275.10 1.451 1.422 – 1.007
CBW-33 200 2 S 10 0.107 166.50 1.449 1.430 – 1.007
The emission of H2 from the soil is large for the Cabauw net-emission experiments, with Fn ranging from 13.7 to 150.2 nmol m−2s−1 and a median value of 41.0 nmol m−2s−1 (Table 2). One experiment, “CBW-28”, shows unusually high emission, with H2increasing to 3010 ppb within 30 min. In comparison, Conrad and Seiler (1980) found aFnof 23–32 nmol m−2s−1for a clover field. Except for the experiments “CBW-28” and “CBW-31”, our Cabauw net-emission experiments are close to theFnfound by them.
The variability inFncould be attributed to different N2fixa- tion flux in our experiments, which could be affected by both spatial density of N2fixation organisms and their N2fixation activities. The N2fixation activity could be regulated by vari- ous factors including temperature, moisture, light availability and carbon storage (Belnap, 2001), which were not measured are therefore not discussed here.
4.2 Fractionation during soil uptake
Fractionation during soil uptake of H2can happen during the diffusion into the soil and due to microbial removal within the soil. To further investigate the factors determiningαsoil, information about the soil cover is provided in Table 1a, b. It
is evident that no large differences exist between the Douglas fir, spruce and beech sites, i.e., the variability between sites is similar to the variability within sites. The small number of experiments impedes examining the possible small dif- ferences between sites. In order to investigate the diffusion effect, we removed the soil cover in experiments “SPU-8”
and “SPU-12” at the same place of experiments “SPU-7”
and “SPU-11”. The removal of leaves (“SPU-8”) and needles (“SPU-12”) increasedαsoilby≈0.014, thus towards smaller fractionation, which indicates that diffusion contributes to the fractionation. Asvd also increases when the soil cover is removed, faster deposition is associated with smaller frac- tionations in these experiments, which is similar to the results from Rice et al. (2011).
Theαsoilfor the Cabauw net-uptake experiments is higher and more scattered than that for the Speuld experiments (0.951±0.026 vs. 0.937±0.008). This could be caused by the interference of D-depleted H2from the strong soil emis- sion in Cabauw, which may not be perfectly captured via the mathematical models applied. As can be seen from the strong decline ofδD with time in the middle panel of Fig. 6, though soil uptake of H2 dominates for the Cabauw net-uptake ex- periments, soil production is still considerable. If part of the
